Direct measurements of NO3 reactivity in and above the boundary layer of a mountaintop site: identification of reactive trace gases and comparison with OH reactivity

NO3 reactivity in and above the boundary layer of a mountaintop site

Direct measurements of NO3 reactivity in and above the boundary layer of a mountaintop site: identification of reactive trace gases and comparison with OH reactivityNO3 reactivity in and above the boundary layer of a mountaintop siteJonathan M. Liebmann et al.

We present direct measurements of the summertime total reactivity of
NO3 towards organic trace gases, kOTGNO3,
at a rural mountain site (988 m a.s.l.) in southern Germany in 2017. The
diel cycle of kOTGNO3 was strongly influenced by
local meteorology with high reactivity observed during the day (values of up
to 0.3 s−1) and values close to the detection limit (0.005 s−1)
at night when the measurement site was in the residual layer and free
troposphere. Daytime values of kOTGNO3 were
sufficiently large that the loss of NO3 due to reaction with
organic trace gases competed with its photolysis and reaction with NO. Within
experimental uncertainty, monoterpenes and isoprene accounted for all of the
measured NO3 reactivity. Averaged over the daylight hours, more
than 25 % of NO3 was removed via reaction with biogenic
volatile organic compounds (BVOCs), implying a significant daytime loss of
NOx and the formation of organic nitrates due to
NO3 chemistry. Ambient NO3 concentrations were measured
on one night and were comparable to those derived from a stationary-state
calculation using measured values of kOTGNO3. We
present and compare the first simultaneous, direct reactivity measurements
for the NO3 and OH radicals. The decoupling of the measurement site
from ground-level emissions resulted in lower reactivity at night for both
radicals, though the correlation between OH and NO3 reactivity was
weak as would be anticipated given their divergent trends in rate constants
with many organic trace gases.

Hydroxyl (OH) and nitrate radicals (NO3) play a centrally important
role in cleansing the atmosphere of trace gas emissions resulting from both
anthropogenic and biogenic activity (Lelieveld et al., 2004, 2016; Ng et al.,
2017). Whereas OH is largely photochemically generated and present at its
highest concentrations during the day, NO3 is generated through the
oxidation of NO2 by O3 and, due to its rapid photolysis
and reaction with NO, is present mainly at night. A further important
difference in the roles of OH and NO3 in the atmosphere is related
to the mechanism of their reactions. NO3 reacts rapidly via
electrophilic addition to unsaturated organic trace gases but reacts
comparatively slowly (via H abstraction) with saturated organics. In the
presence of O2, the initial addition step results in the formation
of nitrooxyalkyl peroxy radicals, which can react with HO2, NO,
NO2, or NO3 to form multifunctional peroxides and organic
nitrates (Fry et al., 2014; Ng et al., 2017).

OH can react both by addition and H abstraction to organic and inorganic
trace gases and may be considered to be more reactive and much less
selective than the NO3 radical. The distinct reaction modes lead to
significant differences in the lifetimes of both radicals, which for OH are
typically less than 1 s and for NO3 can exceed 1 h (Wayne et al.,
1991; Atkinson, 2000; Atkinson and Arey, 2003b; Brown and Stutz, 2012;
Liebmann et al., 2018). Maximum daytime concentrations of OH are typically
less than 1 pptv, whereas NO3 has been observed at the tens to hundreds of
pptv level during night-time (Noxon et al., 1978; Sobanski et al., 2016;
Ng et al., 2017).

The organic nitrates formed in the multistep Reaction (R6) can transfer to
the particle phase or be lost through deposition; N2O5 formed in
Reaction (R4) can react with aqueous particles to form particulate nitrate
and/or ClNO2 (Reaction R7) (Osthoff et al., 2008; Phillips et al.,
2012, 2016; Bannan et al., 2015), thus reducing the rate of photochemical
O3 production (Dentener and Crutzen, 1993). The absolute and
relative fluxes through Reactions (R6) and (R7) thus control to some extent
the lifetime of NOx.

Direct NO3 reactivity measurements have recently become possible
(Liebmann et al., 2017) and the first deployment in a forested region
revealed a large NO3 reactivity at canopy height, not all of which
could be accounted for by simultaneous measurements of a large suite of
organic trace gases (Liebmann et al., 2018). The authors concluded that unmeasured
monoterpenes as well as sesquiterpenes were likely to be responsible. The
difference between the observed (or derived) reactivity and that calculated
from summing loss rates for a set of reactive trace gases is generally termed
“missing reactivity” as frequently reported for OH (Nölscher et al.,
2012). In this paper we prefer the term “unassigned reactivity”.

Previous work on NO3 reactivity has also revealed a strong
meteorological influence on the NO3 lifetime, especially when air
masses are decoupled from the surface layer into which reactive trace gases
(NO and BVOC) are emitted at night (Brown et al., 2007b,
2011; Sobanski et al., 2016; Liebmann et al., 2018).

In this paper we describe direct measurements of NO3 reactivity in
ambient air on a rural mountain site in southern Germany and interpret the
data based on measured VOCs and in terms of the underlying meteorological
situation. We also compare NO3 reactivity to simultaneous measurements
of OH reactivity over the same period.

During the period 20 July to 6 August 2017 NO3 reactivity
measurements were conducted in parallel with ongoing observations at the
Meteorological Observatory Hohenpeissenberg (MOHp) in Bavaria, southern
Germany. The observatory is a meteorological monitoring and Global Atmosphere
Watch site operated by the German Meteorological Service (DWD). It is
located on the Hohenpeissen mountain (988 m a.s.l.), 300–400 m above
the surrounding countryside about 40 km from the northern rim of the Alps
and has been the location of several intensive field campaigns
(Plass-Dülmer et al., 2002; Birmili et al., 2003; Handisides et al., 2003; Mannschreck et
al., 2004; Bartenbach et al., 2007; Hock et al., 2008; Novelli et al., 2017).
The vegetation around the measurement site consists of coniferous trees and
beeches growing on the slopes of the mountain, while grassland and marshes are
dominant in the valley. Tourism-related vehicular emissions represent a
potential source of local anthropogenic pollution, especially at the weekends.
The nearest city, Munich, is about 70 km to the north-east.

Trace gases were sampled into the NO3 reactivity and NO2 CRD
instruments through 1 inch outer-diameter PFA tubing (20 m long, operated at
a flow of 40 dm3 min−1; STP) located 1.5 m above the roof directly
next to the VOC inlet. The inlet was circa 3 m of distance from the inlet used
for the other NOx measurements and circa 2 m of distance from the
OH reactivity inlet.

2.1 NO3 reactivity measurements

The NO3 reactivity instrument was operated in a laboratory located
on the third floor of the MOHp station building at the Hohenpeissenberg.
Air samples were drawn at a flow rate of 2900 cm3 (STP) min−1
through a 2 µm membrane filter (Pall Teflon) and 4 m of PFA tubing
(6.35 mm OD) from the centre of the bypass flow (see above), resulting in
a 7.5 s residence time for the transport of air from the sampling point.
During night-time (≈ 19:00–03:50 UTC) ambient air samples were
drawn through a heated glass flask (35 ∘C, residence time 20 s) to
destroy ambient N2O5 and NO3, which would potentially
interfere with the reactivity measurements. Operational details of the
instrument were recently described by Liebmann et al. (2017). NO3
radicals were generated by mixing NO2 and O3 at elevated
pressure (1.5 bar, ≈5 min reaction time) and passing the mixture
through an oven at ≈ 100 ∘C to convert all
N2O5 to NO3 (Reactions R2–R5). The effluent from the
oven was mixed with either zero-air or ambient air in a flow tube
with the temperature set to 21 ∘C to yield typical (initial) NO3
mixing ratios of 40–60 pptv.

After a fixed reaction time, the remaining NO3 was detected by
cavity ring-down spectroscopy (CRDS) at 662 nm. The lower pressure at the
top of the Hohenpeissenberg station (903±8 hPa) meant that the
reaction time was reduced from 10.5 s as previously reported (Liebmann et
al., 2017) to 9.5 s. The measurement cycle was typically 400 s for
synthetic air and 1200 s for ambient air, with intermittent signal zeroing
(every ≈100 s) by addition of NO. The fractional loss of
NO3 in ambient air compared to zero-air was converted to a
reactivity via the numerical simulation of a simple reaction scheme (Liebmann et
al., 2017) using measured amounts of NO, NO2, and O3. The
parameter obtained, kOTGNO3, is an NO3
loss rate constant from which contributions by NO and NO2 have been
removed and thus refers to reactive loss to organic trace gases (OTGs) only.
Throughout the paper, NO3 reactivity and
kOTGNO3 are equivalent terms, with units of
s−1. The upper measurement limit to kOTGNO3
was 45 s−1, achieved by automated, dynamic dilution of the air sample.
The lower limit was 0.005 s−1, defined by the stability of the
NO3 source. Online calibration of the reactivity using an NO
standard was performed every ≈2 h for 10 min. The uncertainty of
the measurement was between 0.015 and 0.205 s−1, depending mainly on
dilution accuracy, NO levels, and the stability of the NO3 source as
described by Liebmann et al. (2017). Since its first description in Liebmann
et al. (2017), the instrument has been extended with a further cavity to
measure mixing ratios of NO2 (see below).

2.2 NO2, NO, and O3 measurements

Since its first deployment, the NO3 reactivity instrument described by
Liebmann et al. (2017, 2018) has been extended with a further cavity to
measure NO2. This is described here for the first time and thus in
detail. The CRDS measurement of NO2 uses a 2500 Hz, square-wave-modulated, 40 mW laser diode
located in a Thorlabs LDM21 housing and
thermally stabilized at 36 ∘C using a Thorlabs ITC-510
Laser Diode Combi Controller to produce light at 405 nm (0.5 nm full-width
at half-maximum). The laser-diode emission is first directed through an
optical isolator (Thorlabs IO-3D-405-PBS), focused by a lens (Thorlabs
C340TMD-A) into the optical fibre (0.22 NA, 50 µm core,
400–2400 nm),
and then collimated (Thorlabs FiberPort collimator PAF-X-7-A) to a beam
diameter of about 6 mm before entering the cavity. Part of the laser
emission was directed to an Ocean Optics spectrograph to continuously
measure the laser emission spectrum.

The NO2 cavity (Teflon-coated glass; DuPont, FEP, TE 9568, length
70 cm, volume 79 cm3) was operated at 30 ∘C at a flow rate of
3000 cm3 (STP) min−1, resulting in a residence time of
approximately 1 s. To remove particles, air was drawn through a
2 µm membrane filter (Pall Teflon) from the centre of the same
high-flow bypass used for the NO3 reactivity measurements. Light
exiting the cavities through the rear mirror was detected by a
photomultiplier (Hamamatsu E717-500), which was screened by a 405 nm
interference filter. The pre-amplified PMT signal was digitized and averaged
with a 10 MHz, 12 bit USB scope (Picoscope 3424), which was triggered at the
laser-modulation frequency of 2500 Hz.

The ring-down constant in the absence of NO2 was obtained by adding
zero-air every 30 points of measurement for approximately 15 s. The
L∕d ratio (the ratio of the distance between the cavity mirrors,
L, and the length of the cavity that is filled by absorber, d) was
determined as described previously (Schuster et al., 2009; Crowley et al.,
2010) and was 1.00±0.03. Inverse decay constants in dry zero-air at
660 Torr were usually between 28 and 31 µs, indicating optical path
lengths of ≈8–9 km. The measurement precision (6 s integration)
was circa 150 pptv. The cavity was not pressure stabilized, leading to a
pressure difference of circa 2.5 Torr when switching from ambient air to
zero measurements. The data were corrected for the change in Rayleigh
scattering resulting from the pressure difference (typically 120 pptv) and
also different relative humidities (typically 60 to 100 pptv) when switching
from ambient to zero-air measurement as described by Thieser et al. (2016).
The laser spectrum was measured every hour and used to calculate an effective
cross section (≈6×10-19 cm2 molecule−1)
using a literature absorption spectrum (Voigt et al., 2002). The overall
uncertainty of the measurement is mainly determined by the uncertainty in the
cross section (6 %). Other contributions are from NO2 formation
(from reaction of NO with O3 in the inlet lines, ≈0.5 %), the correction for humidity and pressure differences
(5 %), and uncertainty in the L∕d ratio (2 %), giving an
estimated uncertainty of 9 %. The detection limit of the instrument can
be estimated from the variability in the zeros and was usually around
150 pptv.

NO2 measurements were made from 20 July to 4 August with breaks
from 27 July to 2 August and from 4 to 6 August due to instrumental problems.
NO2 mixing ratios were corrected for its formation (Reaction R1)
during transport from the rooftop inlet to the cavity (≈7.5 s).

Two commercially available instruments operated permanently at the site also
provided measurements of NO2 and NO. These were a cavity phase-shift
(CAPS) instrument for NO2 measurement and a chemiluminescence device
(CLD) for NO2 and NO. The CAPS (Aerodyne ambient monitor version 2012)
had a detection limit of 315 pptv (3σ in 1 min integration time)
and an uncertainty of 10 % (1 min integration time). The uncertainty of
the CAPS increased during the period of the campaign due to increased
noise, probably caused by an undetected leakage or mechanical instability
between the lens and the mirror of the measurement cell.

The CLD (ECO PHYSICS, model AL 770 pptv) uses chemiluminescence from the
reaction of NO with ozone in combination with a blue-light converter to
convert NO2 to NO. The instrument was routinely calibrated once a
week (10 ppmv ± 5 % NO in N2; Riessner, Germany).
Deviations between two calibrations are typically below 3 %.
Detection limits during the intensive were 11 pptv for NO, 16 pptv for
NO2 (3σ in 1 min integration time), and median
uncertainties are 27 pptv (7 %) for NO and 70 pptv (10 %) for
NO2 (2σ at 1 ppbv in 1 min integration time). Corrections
were applied to take into account NO loss (≈5 %–30 %) and
NO2 formation (≈0 %–12 %, typically ≈2 %) due to further reactions involving ozone in the inlet tubing.

A comparison of the three NO2 measurement instruments is given in
Fig. S1 of the Supplement, which plots the NO2 mixing ratios
(averaged over 60 s) of the CLD and CAPS instruments versus the CRDS. A
least-squares fit (considering errors in both parameters) to the plot of
NO2 (CLD) versus NO2 (CRDS) has a slope of 0.94±0.25 and
an intercept of 0.00±0.04. For the plot of NO2 (CAPS) versus
NO2 (CRDS) comparison we derive a slope of 0.95±0.02 and an
intercept of 0.00±0.02. Within combined uncertainty, the NO2
measurements are thus in agreement. The NO2 mixing ratios used as input
to calculate the NO3 reactivity were taken from the CRDS instrument,
with data gaps filled by CAPS measurements.

Ozone was monitored with a UV absorption instrument (Thermo Environmental
Instruments Inc., model TECO 49C), which is calibrated at regular intervals
with a transfer standard (TECO 49PS). The uncertainty in the ozone mixing
ratio is 1.2 ppbv or 2 % (2σ in 1 h).

2.3 NO3 measurements

For the measurement of ambient NO3, a 10 m length of PFA tubing
(3∕8 inch outer diameter) was installed on the top of the building circa
10 cm from the VOC inlet. A 2 µm pore PTFE filter (47 mm in
diameter, replaced every hour) in a PFA filter holder was located at the end
of the inlet. The tubing was connected to a bypass pump operated at
20 dm3 (STP) min−1 to reduce the residence time. The sample flow
through the cavity was increased to 8 dm3 (STP) min−1 to reduce
the NO3 residence time within the cavity. NO3 mixing
ratios were recorded every 6 s (3600 ring-downs co-added) with zeroing by
titration (NO addition) every 15 data points. The NO3 transmission
through the inlet (67±15 %), filter and filter holder (84±10 %), and cavity (88±10 %) was established post-campaign as
described by Schuster et al. (2009) and used to correct the data. The overall
uncertainty in the NO3 measurements, including uncertainty in the
absorption cross section, was circa 35 %.

2.4 OH reactivity measurements

OH reactivity measurements were conducted using a chemical ionization mass
spectrometer (CIMS) in which OH radicals (generated by the photolysis of
H2O at 184.95 nm) are converted to H2SO4 (Berresheim
et al., 2000; Schlosser et al., 2009). For the derivation of OH reactivity
(ktotalOH), relative OH radical concentrations are
measured at two fixed reaction times and a decay constant is derived assuming
exponential behaviour. After correction for wall losses, as well as
NO-induced HOx recycling in the sample tube, ambient
reactivities between 1 and 40 s−1 are measurable. OH reactivity
measurements were made every 20 min throughout the measurement period.
Measurements were discontinued during periods of precipitation and when the
pinhole to the mass spectrometer vacuum system was blocked by insects. The
instrument performs best if NO mixing ratios are below 4 ppbv and
reactivities do not exceed 15 s−1; for the measurements reported here,
the mean uncertainty in the OH reactivity was ±1.2 s−1 (or
46 %). OH reactivity calibration was carried out before and after the
measurement period, and the calibration factor was applied to the whole
dataset. The determination of the OH wall loss rate from zero reactivity
measurement (null measurement) using synthetic air cylinders was not reliable
and therefore the zero was estimated using night-time measurements when
sampling from above the boundary layer. Details are given in the Supplement.

2.5 VOC measurements

A gas chromatograph (GC-MS/FID model Agilent 6890 with 5975 B inert XL MSD)
was used for the detection of C5–C13 NMHCs and BVOCs (Hoerger et
al., 2015). In a custom-made pre-concentration unit, air was sampled at
30 ∘C on a three-bed adsorption trap and, after a cryo-focussing step,
injected onto the GC column (50 m BPX-5). Subsequently, signals were
detected with a mass spectrometer (MS) running in parallel with a flame
ionization detector (FID). The instrument measured i.a. isoprene and a wide
variety of monoterpenes with uncertainties (2σ) from 6 % to
100 % depending on the compound.

For the detection of light NMHCs (C2–C8), a GC-FID system (GC-1,
Varian 3600 CX, FID detector) described in detail by Plass-Dülmer et
al. (2002) was used. In both systems, an ozone scrubber (impregnated filter
with Na2SO3) was used and water was removed from the sample air
either by hydrophobic adsorbents (C5–C13) or a cold trap
(C2–C8) prior to the pre-concentration step. VOCs were sampled
every hour for either 15 (C2–C8) or 20 min (C5–C13).
During the rainy period from 24 July 12:00 UTC to 27 July 12:00 UTC, VOCs
were only measured twice daily.

2.6 Particle measurements

The aerosol surface area was calculated using particle number size
distributions obtained from a custom-built SMPS described in detail in
Wiedensohler et al. (2012) and Birmili et al. (2016). Briefly, the instrument
uses a Vienna-type DMA with a condensation particle counter (CPC model 3772,
TSI Inc.) to measure particles between 10 and 800 nm. The sheath flow rate
is 5 L min−1 at an aerosol flow rate of 1 L min−1; both are
actively dried. The typical time resolution for one combined up-scan and
down-scan is 5 min.

Figure 1 displays the time series of NO3 reactivity
(kOTGNO3) along with related trace gases and
meteorological data obtained during the intensive. Sunrise was around
03:50 UTC and sunset at ≈19:00 UTC. Two mild days (Tmax=20–25 ∘C) at the beginning of the campaign were followed by a
3-day period with heavy rains and maximum temperatures of around
10 ∘C and then by a warm period with temperatures up to
30 ∘C and occasional thunderstorms. The predominant wind direction
was west–south-west with only minor contributions from other directions
(Fig. S2). Wind speeds were generally around 2.5 to 7.5 m s−1,
increasing up to 15 m s−1 during the rainy periods. The highest values
of kOTGNO3 were detected with north-easterly winds
(Fig. S2) coincident with the warmest days of the campaign and the highest
biogenic emissions (see Sect. 4.2). Ozone mixing ratios were strongly
correlated with temperature and ranged from 85 ppbv on 1 August during the
warm, photochemically intense period to less than 20 ppbv during the cool,
rainy period between 23 and 28 August. NOx levels
(NOx=NO+NO2) during the intensive
were generally between about 0.5 and 4 ppbv. The mixing ratios of NO, a
trace gas which can potentially impact NO3 lifetimes, were
generally below the detection limit (≈12 pptv) during most of the
nights, increasing to maximum values of <1 ppbv during the day.
Occasional maxima of more than 1 ppbv NO were observed due to local traffic.

3.1 NO3 reactivity

NO3 reactivity was measured continuously during a 3-week
intensive (20 July to 6 August 2017) with the exception of one night
(2–3 August) when, using the same instrument, NO3 mixing ratios
were measured instead. The full dataset of kOTGNO3
is reproduced in Fig. S3 of the Supplement together with the corresponding
95 % uncertainty limits, which take into account drifts in the zero
signal, the stability of the NO3 source, uncertainty in the
dilution factors, uncertainty of the NO and NO2 mixing ratios, and
the corresponding rate constants.

Figure 2(a)kOTGNO3 (black) and O3
mixing ratios (red) from 29 to 30 July. From 23:50 UTC until sunrise
the measurement site is located in the residual layer and free troposphere.
(b) Temperature (T), relative humidity (RH), and wind direction
(WD) during the same period.

As described above, during daytime the short NO3 lifetime normally
results in levels that are under the detection limit of most instruments,
precluding the estimation of NO3 reactivity via stationary-state
calculations based on its mixing ratio and production rate. In contrast, our
direct measurement enables us to derive the NO3 reactivity over the
full diel cycle. During the intensive, the 10 min averaged values of
kOTGNO3 ranged from below the detection
limit (<0.005 s−1) to values as high as 0.3 s−1. Campaign-averaged values were low (≈0.01 s−1) during night-time but a
factor of 10 larger ≈0.1 s−1 at 14:00 UTC (local 16:00) and
more variable during daytime.

This observation is in stark contrast to the high relative night-time–daytime
NO3 reactivities we observed in a boreal forest (Liebmann et al.,
2018) and is related to very different meteorological conditions at the two
sites. In the boreal forest, the canopy-level NO3 reactivity was
controlled by the rate of emission of biogenic VOCs into a nocturnal boundary
layer of varying height and stability. The elevated location of the
Hohenpeissenberg observatory, located on a mountaintop above the surrounding
countryside, favoured sampling from the residual layer and free troposphere at
night-time. In the absence of turbulent exchange, the residual layer and free
troposphere may become disconnected from the planetary boundary layer (PBL)
and thus from ground-level emissions of reactive trace gases; the layers may thus
contain low levels of biogenic trace gases as well as low(er) levels of
NO2 and higher levels of ozone (Aliwell and Jones, 1998; Allan et
al., 2002; von Friedeburg et al., 2002; Stutz et al., 2004; Brown et al.,
2007a, b; Brown and Stutz, 2012). NO3 lifetimes as long as 1 h
(using stationary-state analyses) have been reported for mountain sites when
sampling air from above the nocturnal boundary layer (Brown et al., 2016;
Sobanski et al., 2016).

During the Hohenpeissenberg intensive two distinct air mass types were
encountered at night, whereby values of kOTGNO3
were either at (or below) the detection limit or well above it (referred to
as “type 1” and “type 2”, respectively). Figure 2 displays a time series
of kOTGNO3 over a single night (29–30 July) in
which a switch from type 2 to type 1 was observed. From early evening until
shortly before 12:00 UTC, NO3 reactivity was variable with values
between circa 0.02 and 0.03 s−1. A sharp reduction in kOTGNO3 was then observed with values close to the detection limit
until sunrise (04:00 UTC). The reduction in kOTGNO3 was accompanied by a drop in relative humidity (from
70 % to 60 %) and an increase in O3 (46 to 52 ppbv), both
clear indicators of sampling from the residual layer. At the same time, the
wind speed increased and the temperature became more variable, indicating
that the site was close to the inversion level. At first sunlight, turbulent
mixing resulted in gradual connection of the boundary layer and overlying
residual layer, leading to an increase in kOTGNO3.
Upslope winds caused by heating of the mountainside may also have enhanced
transport of air masses with high reactivity to the measurement site. Median
diel profiles of NO3 reactivity on type 1 (altogether five) and
type 2 nights (altogether 10) are shown in Fig. 3. With the exception of the
very low reactivity during type 1 nights, type 1 and type 2 have similar diel
shapes and similar maximum reactivities.

3.2 NO3 reactivity calculated from VOC measurements

In this section we assess the contribution of various VOCs to the observed
NO3 reactivity. The most abundant BVOCs were isoprene, sabinene,
α-pinene, and β-pinene with maximum mixing ratios during the
warm period around August. The time series of BVOC mixing ratios are
displayed in Fig. S4 of the Supplement. kOTGNO3 is
a total loss rate constant for chemical reactions of [NO3] with all
organic trace gases present and can be compared to the summed loss rate
constant (kVOCNO3) (also in units of s−1)
obtained from the concentrations of individual VOCs in the same air mass,
[Ci], and the rate coefficient (ki):

(1)kVOCNO3=∑kiNO3[Ci],

where [Ci] is the measured BVOC concentration and ki the
corresponding rate constant. Individual values of kVOCNO3, calculated using rate constants from the IUPAC evaluation
(IUPAC, 2018) or elsewhere in the literature (Shorees et al., 1991), are
plotted with interpolated 20 min averages of kOTGNO3 as a time series in Fig. 4a.

The data are also displayed as a pie chart in Fig. 4b in
which the contribution of individual biogenic trace gases to the
NO3 reactivity are listed. Of the terpenoids, α-pinene
contributed most to the overall NO3 reactivity (≈16 %)
followed by sabinene (≈12 %) with other individual BVOCs
contributing less than 10 %. VOCs such as methanol, acetaldehyde, ethanol,
acetone, methylethylketone, alkanes, and aromatics were also measured but not
included in calculations of
kVOCNO3 as their summed
contribution reached a maximum of 1.5×10-4 s−1 and was on
average 5×10-5 s−1. As kVOCNO3 and kOTGNO3 show a
similar dependence on wind direction (Fig. S2) and because only BVOCs were
used for the derivation of kVOCNO3, we
conclude that the high NO3 reactivities measured in air masses arriving
from the east and north-east are mostly from trace gases of biogenic origin.

Figure 4(a) Measured values of kOTGNO3 (black)
compared to the loss rate constant attributed to individual VOCs. The term
“other” includes terpinolene, β-phellandrene, α-terpinene,
γ-terpinene, α-thujene, and camphene. Myrcene and
α-phellandrene were also measured but below the detection limit during
the whole campaign. The pie chart (panel b) indicates the campaign-averaged contribution of each measured VOC to the NO3 loss rate as
well as reactivity that could not be attributed to measured VOCs
(“unassigned reactivity”).

Type 1 nights were characterized by very low BVOC mixing ratios, sometimes
below the detection limit, whereas isoprene was still present. This
observation is consistent with a long lifetime for isoprene in the residual
layer at night (Brown et al., 2007a) as the OH concentration is too low
and the NO3 reaction too slow to remove it efficiently. Under
conditions of very low NO3 reactivity, the fractional contribution of
isoprene to the overall reactivity could increase to ≈100 %
(from typically 20 % during the day). During type 2 nights (those with
non-zero NO3 reactivity) isoprene and monoterpenes were always detected
and monoterpenes were the dominant reaction partners for NO3.

The difference between kOTGNO3 and
kVOCNO3 (i.e. the NO3 reactivity not
accounted for by measured VOCs) is defined as unassigned reactivity
(s−1).

(2)unassignedNO3reactivity=kOTGNO3-kVOCNO3

A plot of kVOCNO3 versus kOTGNO3 (see Fig. S5 of the Supplement) has a slope of 1.55 and an
intercept of 0.005. This implies, on average, an unassigned reactivity of
≈34 % when kOTGNO3=0.3 s−1 and
a unassigned reactivity of ≈50 % when kOTGNO3=0.03 s−1. However, both kOTGNO3 and kVOCNO3 are associated with
uncertainty, which needs to be rigorously assessed to test whether the
unassigned reactivity is significant. To do this we propagated uncertainty in
each of the terms ∑kiNO3[Ci]
(mainly related to VOC measurements and assuming 15 % uncertainty in the
rate coefficients) and derived mean diel profiles of kOTGNO3 and kVOCNO3 for the whole campaign
(hour averages). We note that IUPAC-listed uncertainties for each VOC are
generally larger than 15 % (especially when only few studies are
available) and the use thereof would substantially increase the uncertainty
in kVOCNO3. The results are shown as a time series
and a campaign-averaged diel profile in Fig. 5. For the time series, we plot
data points only when values of both kVOCNO3 and
kOTGNO3 were available. The time series illustrates
that, within combined uncertainty, the data overlap so that unassigned
reactivity is not significantly distinct from zero apart from a few isolated
data points. This contrasts with the conclusions of a previous campaign (boreal
forest) with this instrument (Liebmann et al., 2018) in which up to
40 %–60 % of the measured reactivity could not be accounted for by
measured VOCs.

Figure 5(a) Time series of kVOCNO3 and
kOTGNO3. The error bars on the
kOTGNO3 measurements are total uncertainty calculated as
described by Liebmann et al. (2017). The uncertainty in
kVOCNO3 (shaded red area) is dominated by uncertainty in
the mixing ratios of the VOCs. Panel (b) is a campaign-averaged diel
profile of the same parameters, whereby the uncertainties also include
variability.

The comparison of measured NO3 reactivity with that calculated from VOC
measurements suffers from substantial uncertainties in both parameters. For
this campaign, the uncertainty in NO3 reactivity was larger than
previously reported (Liebmann et al., 2018) due
to reduced stability of the NO3 source, which will be improved in
future versions of the instrument. A further obstacle to the calculation of
unassigned reactivity was the uncertainty in the measurements of biogenic
VOCs and the associated rate constants for the NO3 reaction. The latter
could be reduced by more plentiful and more accurate data on NO3 rate coefficients with
some of the more important biogenic species.

Although rough estimates of NO3 concentrations at the
Hohenpeissenberg have been made (Handisides et al., 2003; Bartenbach et al.,
2007), no direct NO3 measurement had been previously made. For this
reason, on just one night during the intensive (2–3 August), the instrument
was modified to enable measurement of ambient NO3 mixing ratios
rather than NO3 reactivity. The NO3, O3, and
NO2 mixing ratios and meteorological data are plotted in Fig. 6.

NO3 mixing ratios slowly increased in the first half of the night,
reaching a maximum of 13 pptv around 21:40 UTC. At this time the
O3 mixing ratios were also the largest and highly variable. After
≈22:30, O3 was slowly removed, and the NO3 decreased
by a factor of 10 or more, indicating that the instrument was sampling more
reactive boundary layer air. This is also evident in the increase in
relative humidity and decrease in the temperature until about 01:30.

Given sufficient time, stationary state can be reached for NO3 at night
in which the production and loss terms are approximately balanced (Brown
et al., 2003a; Crowley et al., 2010, 2011). In this case
NO3 mixing ratios can be described by the ratio of their production
rate and loss rate (Eq. 3).

(3)NO3ss=NO3productionrateNO3lossrate

The production rate is governed by the [NO2] and [O3] mixing ratio
and the corresponding rate constant k2. If the loss processes are due to
reaction with VOCs only, this expression becomes

(4)NO3ss=O3NO2k2kVOCNO3.

During this night kOTGNO3 was not measured
so kVOCNO3 was used to account for
NO3 losses.

Figure 7 shows the measured NO3 mixing ratios (black) compared to those
derived from Eq. (4) using the measured VOC concentrations (red curve).
Clearly, the predicted stationary-state NO3 concentrations are too
high (by a factor of up to 3–4), implying that other NO3 loss processes
must be considered. As the directly measured reactivity, kOTGNO3, agrees rather well with that derived from VOC
measurements (kVOCNO3) calculated on other
campaign nights, it would seem unlikely that unmeasured VOCs contribute
sufficiently to NO3 losses to explain this large factor.
Stationary-state concentrations of NO3 are influenced not only by VOCs
but also by NO (if present at night) and also indirectly via the heterogeneous
loss of N2O5. Equation (5) can be extended to include these
processes (Martinez et al., 2000; Geyer et al., 2001; Brown et al.,
2003a, b, 2009; Crowley et al., 2010;
Sobanski et al., 2016).

(5)NO3ss=O3NO2k1kVOCNO3+k2NO+K5NO2fhet,

where K5 is the equilibrium constant for the forward and reverse
Reactions (R4) and (R5). The loss frequency due to the heterogeneous uptake of
N2O5 to particles (fhet) can be calculated by
Eq. (6):

(6)fhet≈γc‾A4,

which is approximately valid if the particles are less than ≈1µm in diameter. In this expression, A is the aerosol surface
area density (cm2 cm−3), c‾ is the mean molecular
velocity of N2O5 (26 233 cm s−1 at 298 K), and γ
is the dimensionless uptake coefficient. If we assume a large value for the
uptake coefficient of 0.03 as characteristic for aerosol with low organic
content (Bertram and Thornton, 2009; Bertram et al., 2009; Crowley et al.,
2011; Phillips et al., 2016) and use the aerosol surface area density of
1.25–1.55×10-6 cm2 cm−3 (measured by a scanning
mobility particle sizer for 10–890 nm), we obtain values for
fhet of 2.4–2.9×10-4 s−1. Incorporation of
the heterogeneous loss of N2O5 reduces the predicted NO3
by only about 5 pptv as shown in Fig. 7 and cannot reproduce the observed
NO3 mixing ratios. The efficient reaction of NO with NO3
means that low mixing ratios of NO can contribute to NO3
reactivity, and we calculate that ≈5 pptv NO would approximately
align the calculated NO3 mixing ratio with that measured for much
of the night. This value is, however, below the detection limit of the CLD
(11 pptv) used to measure NO and we cannot conclude that NO at such levels
was responsible for the reduction in NO3 levels required to bring
observations and steady-state calculations into agreement.

The discussion above demonstrates that the calculation of
NO3 reactivity from stationary-state calculations can be precarious and
subject to large cumulative uncertainty from e.g. measurement uncertainty in
NO3 mixing ratios, uptake coefficients, aerosol surface area, and NO
mixing ratios close to instrumental detection limits.

Figure 8(a) Stationary-state NO3 mixing ratios calculated using
kOTGNO3, [NO]k3, K5[NO2]fhet,
and JNO3 for the entire campaign and comparison with the
measured NO3 mixing ratios (3 August). Panel (b) plots the
time series of production and loss rates used for the calculation of
[NO3]ss.

Figure 8 shows the NO3 production rate (panel b, black curve)
and total loss rate (panel b, red curve) as well as the stationary-state
NO3 mixing ratios for the entire intensive period (Fig. 8a, black curve). During nights on which the reactivity fell below the
detection limit of the instrument kOTGNO3 was set
to 0.005 s−1. Calculated NO3 mixing ratios were in the
sub-pptv range during daytime and around 1–15 pptv during night-time. The
NO3 mixing ratios thus derived are comparable to those measured on
a single night (Fig. 8, red curve) and are broadly consistent with previous
estimates for this site (Handisides et al., 2003; Bartenbach et al., 2007).

3.4 Contribution of NO3 reactivity to NOx loss

At night-time, in the absence of NO and sunlight, each NO3 radical
formed in the reaction of NO2 with O3 will either be
removed indirectly via the uptake of N2O5 onto particles or will
react with a biogenic hydrocarbon. The latter results in the formation of an
organic nitrate at a yield of between 20 and 100 %, depending on the
specific VOC (Ng et al., 2017) and hence the removal of NOx.
The large daytime values for kOTGNO3 obtained in
this study suggest that even during sunlight hours (when NO3 is
generally considered to be of little importance owing to its rapid photolysis) significant amounts of
NO3 form organic nitrates rather than reforming NO2 by
reaction with NO or photolysis.

The fraction, f, of NO3 that will react with organic trace gases is
given by

(8)f=kOTGNO3kOTGNO3+JNO3+NOk3+K5NO2fhet,

where the denominator sums all loss processes for NO3. Mixing
ratios of NO below the 11 pptv detection limit were set to zero. Figure 9
illustrates this via a diel cycle of the median for f. At night-time,
≈99 % of the NO3 will be lost to reaction with BVOCs,
with indirect heterogeneous losses representing the remaining 1 %. Note
that the presence of NO at 5–10 pptv levels during the night would
significantly reduce this value (see Sect. 3.3).

Figure 9The fraction, f, of the total NO3 loss with organic trace
gases as a campaign mean diel cycle where f=kOTGNO3/(kOTGNO3+JNO3+[NO]k3+K5[NO2]fhet). The error bars reflect variability only and
do not consider systematic uncertainty.

During daytime, at the peak of the actinic flux (max JNO3≈0.2 s−1) and correspondingly high levels of NO (kNO=0.1–0.2 s−1), 20 % of the formed NO3 was lost due to
reaction with organic trace gases, increasing up to 40 % in the late
afternoon. This result is comparable with reactivity measurements in a boreal
forest in Finland during IBAIRN 2016 in which a very similar diel profile for
f was determined (Liebmann et al., 2018). The NO3 reactivity data
from these measurements indicate that the role of NO3 as a daytime
oxidant of biogenic VOCs in forested regions may so far have been
underestimated, which in turn has implications for understanding the diel
cycle of organic nitrate and secondary organic aerosol formation in such
environments.

As mentioned above NO3 radicals and OH radicals react with atmospheric
trace gases via different mechanisms, resulting in profoundly different rate
coefficients and thus reactivities. By combining continuous, on-site
measurements of OH reactivity with the NO3 reactivity measurements
during the intensive period, we were able to generate the first dataset of
the simultaneous, direct measurement of OH reactivity,
ktotalOH, and NO3 reactivity at any location.

Figure 11Correlation between OH reactivity and NO3 reactivity. The
coloured lines are relative NO3 and OH reactivity for single VOCs.
The measured NO3 and OH reactivities are depicted as black (night-time) and
orange data points (daytime).

To aid comparison, we subtracted the contributions of several inorganic and
organic trace gases (NO, NO2, SO2, CO, CH4)
that are not included in kOTGNO3 or do not react to
a significant extent with NO3 from the total OH reactivity and thus
derived kOTGOH.

kOTGOH=ktotalOH-kNOOHNO-kNO2OHNO2-kSO2OHSO2(9)-kCH4OHCH4-kCOOHCO

Figure 10 depicts the time series of kTotalOH,
kOTGOH, and kOTGNO3. All
display maximum values close to midday, though
kOTGOH values averaged over the intensive are larger by a
factor of 44 than kOTGNO3, reflecting
generally larger rate coefficients for OH. The blue shaded areas for
kTotalOH represent the 1σ uncertainty of
the measurements. Total uncertainty in kOTGOH and
kOTGNO3 is not shown to preserve clarity of
presentation. The time series of kOTGOH can be
found in the Supplement (Fig. S6) where it is compared to that calculated
from individual VOCs at the corresponding rate coefficient,
kVOCOH. As is often the case for OH, there are periods
in which the calculated reactivity significantly underestimates (up to a
factor of ≈4) the measured values. A detailed comparison of measured
and calculated OH reactivity at this location and including data over a much
longer time period will be the subject of future publications.

The measured reactivities of both radicals show a clear diel profile, with
higher daytime and lower night-time values. Figure 11 shows a correlation plot
of OH and NO3 reactivity divided into day (red) and night-time
(black) data. During the day, the data are highly scattered, which can be
understood when one considers the highly variable organic content of the air
masses being sampled. To illustrate this, we have drawn the expected
correlation lines (based on the known relative rate coefficients) for single-component organic trace gases including isoprene and terpenes. The expected
slopes for these individual VOCs are very different and encompass the full
scatter in the observations, which is the result of changing atmospheric
composition (i.e. the mix of reactive organic species) owing to changes in
air mass age and source region (wind direction) during the campaign. The
extremes are represented by α-terpinene (which favours NO3)
and CH4 (which favours OH).

During night-time (black points) the plot of kOTGOH
versus kOTGNO3 is less scattered, indicating that
the air masses sampled (often from the residual layer) are chemically less
complex and variable. The data lay close to the line that marks the expected
correlation if isoprene were the dominant sink of both NO3 and OH
at night-time once molecules such as CO and CH4 have been removed
from the term describing OH reactivity. This is in broad agreement with our
observation that isoprene is the main sink of NO3 during nights
when the measurement site was decoupled from direct boundary layer emissions.

Direct measurements of the NO3 reactivity towards organic trace
gases, kOTGNO3, were conducted at the top of the
Hohenpeissenberg (988 m a.s.l.) during an intensive measurement
campaign in the summer of 2017. NO3 reactivities had a distinct
diel profile with values as large as 0.3 s−1 during daytime but close
to or below the detection limit of the instrument during night-time when the
measurement site was frequently in the residual layer and free troposphere.
Within experimental uncertainty, the high daytime NO3 reactivity
was accounted for by BVOCs that were measured at the site and was dominated
by monoterpenes, especially α-pinene and sabinene. On average, and
assuming NO levels below the 11 ppt detection limit to be zero, the reaction
with VOCs accounted for ≈99 % of the loss of NO3
during night-time and an average of 20 % at noon, increasing to
30 %–50 % during early morning and late evening. The reaction of
NO3 with BVOCS therefore represents a significant
NOx loss not only during the night but also during daytime
and implies significant formation of organic nitrates via NO3
reactions throughout the diel cycle. Stationary-state daytime and night-time
NO3 mixing ratios were calculated using the production term and
kOTGNO3 and were broadly consistent with direct
measurement made on one night. A comparison between directly measured OH and
NO3 reactivities was performed, indicating a weak correlation
during the day when chemically reactive, complex, and variable air masses were
encountered. A tighter correlation, consistent with isoprene dominating the
(low) NO3 reactivities, was observed at night.

We would like to thank the DWD for hosting and supporting this measurement
campaign. We would like to thank the DWD personnel for the data evaluation,
most notably Katja Michl for NMHCs, Jennifer Englert for OVOCs, and
Harald Flentje for the SMPS data as well as the great technical support; we
are particularly grateful to Erasmus Tensing, Thomas Elste, and Georg Stange.
We thank Chemours for provision of the FEP sample used to coat the
CRD cavities.

The article processing charges
for this open-access publication were covered by the Max
Planck Society.

We present direct measurements of the summertime total reactivity (inverse lifetime) of NO3 towards organic trace gases at a rural mountain site. High daytime and low night-time values were found. The reactivity was dominated by reaction with monoterpenes and sufficiently high to compete with photolysis and reaction with NO during daytime. NO3 radical measurements from one night are presented. For the first time, direct measurements of OH and NO3 reactivity are compared.

We present direct measurements of the summertime total reactivity (inverse lifetime) of NO3 ...